Abstract Soil organic carbon (SOC) plays a vital role in soil formation. The accumulation of SOC is one of the initial soil forming processes and is determined by physical, chemical, biological and anthropogenic factors with complex interac-tions. On the other hand, SOC and its composition influences other soil forming processes like leaching of cations, soil acidification, gleying including Fe-reduction and podzolization. As SOC is strongly correlated with soil organic nitrogen (SON) and nitrogen being the most widespread constraint for biomass production on cropland, SOC content and composition is a determining factor for soil productivity on well drained soils. Thus, SOC is an effective contributor to the supporting ecosystem services of soil formation on the global land surface and at the same time it positively affects the provisioning ecosystem services (ESs) for supplying food, feed and fiber.

Soil formation is an environmental process which transforms the geological sub-strate (the so-called parent material) into soil through the combined influence of physical, chemical, biological and anthropogenic factors. Jenny (1941) classified the factors of soil formation into the five categories (i) climate, (ii) parent material, (iii) topography, (iv) organisms (including humans) and (v) time (Fig. 17.1).

Soil formation

Parent material

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Fig. 17.1 Five determining factors for soil formation (arrows denote the major direction of influence) (adapted from Jenny 1941)

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These factors drive many different soil forming processes which determine the development of soil properties over time. Thus, soil formation is the result of dynamic processes though at longer time scales (from decades to millennia). Accumulation of soil organic carbon (SOC) is among the initial soil forming processes and is, as such, subject to the factors of soil formation. For example, the organisms living in and growing on a soil determine, in interaction with climatic conditions and time, the amount of carbon (C) which is transferred from the atmo-sphere through C assimilation (i.e., photosynthesis), into the soil via root exudation, root residues and surface litter. At the same time, the amount and quality of SOC influences other soil forming processes.

17.2 Accumulation of Organic Carbon in Soils

The initial state of soil formation is characterized by the decomposition of the parent material through physical and chemical weathering. This process leads to the break-down of larger mineral structures into smaller and less solid fragments. A distinct pore system develops between these fragments which offers space for organisms including the roots of higher plants. If the parent material consists of non-consolidated material, a pore system is already in place, which allows colonization by soil biota and roots. As the organisms on and in the soil die, their biomass, which is composed of about 50 % C bound in organic components, remains in the soil and is subject to various chemical transformations. These include oxidation (in the pres-ence of sufficient oxygen) or fermentation (in the absence of oxygen) catalyzed by different soil organisms. Roots of higher plants release exudates consisting of organic compounds which are easily decomposable. In all cases, some portion of the organic C is released as carbon dioxide (CO2) to the atmosphere, whereas the remaining organic C is either incorporated into the biomass of the organisms or, in the case of some larger molecular organic compounds, may accumulate in the soil. Under natural conditions, accumulation of SOC is the result of an excess C input compared to C output, the C output consisting mainly of CO2 as a product of oxida-tion of organic compounds metabolized by the soil biota. There are three dominant factors which cause a reduction of the C output and hence result in an accumulation of SOC. In mineral soils which are aerated for the major part of the year, the decom-position of organic compounds is reduced by physical or chemical protection against the decomposing activity of soil organisms (Von Ltzow et al. 2006). Also, extremely high soil acidity may cause a reduction in decomposition rate. In organic soils which are saturated with water for more than 300 days per year, the decompo-sition of organic compounds is reduced by oxygen deficiency.

Organic soils which are almost permanently saturated with water are showing soil layers with high concentrations of organic C. Those soil horizons containing at least 12 % (w/w) of organic C are classified as histic horizons (H-horizon). In many cases the amount of organic C exceeds 30 % and organic layers may extend to a thickness of several meters. According to the World Reference Base for Soil

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Resources, soils with an H-horizon are classified as Histosols (Fig. 17.2). They occur usually in the humid regions of the world where an excess of water in the soil is favored by higher annual precipitation compared to potential evapotrans-piration. However, saturation of the soil can also be caused by shallow groundwater or stagnating water, which is the case in e.g. the permafrost zone, where water from the melting snow cannot percolate through the soil profile due to the impermeable permafrost layer in the subsoil.

Generally, in mineral soils, organic C accumulation is highest in the topsoil due to a denser root system and due to litter fall which, in the first instance, is incorporated into the surface soil. Thus, with increasing accumulation of SOC within a mineral matrix over time, a soil layer develops which is called the A-Horizon and which is distinctly different from the layers below due to its dark-grey to blackish color. Shallow soils (soil depth below 30 cm) that consist exclusively of an A-Horizon with underlying hard rock are classified as Leptosols (WRB 2006, Fig. 17.2). Leptosols are typical soils at the initial state of soil formation. As soil weathering proceeds to deeper layers or on non-consolidated parent material, the accumulation of organic C in the soil may continue. In continental climates with hot, dry summers and extremely cold winters, bioturbation is observed. The latter describes the incorporation of organic compounds into the subsoil through the activity of larger soil organisms (macro-fauna). Bioturbation favors the extension of the A-horizon into deeper soil layers forming dark-colored, humus-rich, well-structured soils which are classified either as Chernozems, Kastanozems or Phaeozems (WRB 2006, Fig. 17.2).

Human activities have brought along other types of SOC accumulation in mineral soils, which resulted from the application of large amounts of charcoal and ashes or organic and mineral material from livestock. The former case has

been observed in the Amazon basin where indigenous practices of soil fertility restoration have promoted the formation of so-called Terra preta soils. During medi-eval times, farmers in some regions of Northern Europe used to collect litter material and mineral soil in the forest, brought it home and mixed it with animal manure. Then, this mixture was applied in large quantities to the cropland leading to elevation of the plots and forming so-called Plaggen soils. Both soil types are known to be more fertile and productive compared to the associated soil types which have developed under natural conditions.

In peat bog areas under permanent water saturation, organic C accumulation may proceed over longer time scales, forming deep soils consisting almost exclusively of organic matter (OM). There, organic C accumulation slows down when the water saturation is reduced either by changing climatic conditions or when the H-horizon is elevated above the level of the ground or perched water table. In mineral soils, depending on climate, soil texture, groundwater influence and vegetation, SOC accumulation reaches steady-state conditions with an equilibrium between C input and output. Stable boundary conditions over a longer period of time, result in an equilibrium level of SOC in mineral soils. However, as soon as one of the boundary conditions changes, the equilibrium is disturbed and the SOC concentration will shift to a new equilibrium level. Due to the relatively fast break-down of organic compounds compared to their synthesis, reaching a higher equilibrium concentration requires more time than dropping to lower concentration (Fig. 17.3).

As soon as organisms start to colonize the soil the organisms as well as the accumu-lating organic C interact with various processes of soil formation. Polysaccharides produced by roots or by the excreta of earthworms promote soil structure formation, i.e., the aggregation of individual mineral particles. In addition, high-molecular humic acids create bounds with primary or secondary clay minerals to form stable organo-mineral complexes.

The breakdown of OM by soil biota produces large amounts of CO2 in the soil which transforms into carbonic acid under the presence of water filled soil pores. Subsequently, carbonic acid causes a decrease in pH and an increase in proton concentration in the soil solution. The protons are in equilibrium with other cations like Ca2+, Mg2+ and K+ at the cation exchange sites of the soil matrix. The change in proton concentration leads to desorption of the cations from the soil matrix into the soil solution which makes them prone to leaching. Thus, SOC and its transformation processes promote soil acidification and leaching of neutral cations.

In contrast to soil acidification caused by organic acids and leaching we observe processes which generally buffer acids. There are two simple rules in the soil ecosys-tem. First, all oxidizing processes produce H+ and cause acidification. Second, all reduction processes consume H+ and raise the pH. How does it work? The oxygen dissolved in soil solution is used for respiration by roots, soil macro- and mesofauna and especially by microorganisms. Under water saturated conditions, respiration decreases the oxygen concentration in the soil and due to slow diffusion of oxygen in the soil water, diffusion is not able to replace the consumed oxygen at a sufficient rate. When almost all oxygen is used for respiration, specialists can use chemically bound oxygen for anaerobic respiration. To facilitate these processes, easily available organic soil material must be present. Then, SOC is oxidized and the oxidant is reduced as shown in the following equation using the example of iron oxides (Fe2O3):

C H O Fe O CO H O Fe6 12 6 2 3 2 224 6 6 8+ + + +

The organic component could be sugar or other water soluble organic sub-stances. As shown in the formula above, all these processes produce CO2 and water. They also produce energy. This is the main principle for all aerobic or anaerobic respirative processes. Stable organic molecules like cellulose or lignin will not be mineralized under these conditions, but rather accumulate in the soil. Therefore, under those conditions, OM and thus SOC is enriched in soils. Quite often this is involving material belonging to the group of highly polymerized organic compounds with phenolic ring structures produced from organic matter transformation, but these compounds can also originate from physically or chemi-cally protected litter components. When oxygen levels are too low, we observe a certain chain of substances (ions) which can be reduced by anaerobic reduction, like NO3, Mn4+, Fe3+, SO42, CO2, and H2O. The corresponding products which will be reduced by the reduction processes are N2O, N2, Mn2+, Fe2+, S2, CH4, and

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H2 (Stahr et al. 2008). There are many more ions, which can be involved in the reduction and oxidization chains, like Cu, Co or Mo. They generally do not play a significant role because of low concentrations or contents. With reducing and oxidizing processes, several other products, besides OM, are accumulated. This is iron oxide together with manganese oxide within a fluctuating ground water table (gleying conditions), but also sulfides can accumulate, particularly under perma-nent reduction. Beside these reaction products, other substances of high climatic relevance can be formed like N2O and CH4. All these soil processes can occur in different soil depths. Generally, the processes decelerate with increasing soil depth, because they need special easily available organic C sources. Soils which are temporarily saturated with water either by the influence of ground water (gleyic) or surface water (stagnic) can be very productive with favorable tempera-ture regime and nutrient supply. Then they also tend to be rich in SOC. On the other hand, when these soils are poor, they tend to accumulate less SOC.

Podsolization is a process which cannot occur without contribution of SOC. However, it is completely different from the above described aerobic and anaerobic processes. Podzolization occurs under humid to per-humid climates in soils with free drainage and unbuffered conditions. While keeping an eye on the occurrence of those soils which are characterized by a spodic horizon being enriched in Fe, Al hydroxides and/ or SOC, we generally find these soils under either of the three following condi-tions: These are, first the boreal coniferous forests, secondly the alpine tree line, and thirdly the Fe-poor and extremely sandy soils in other humid climates. The podzoliza-tion process is one of the early soil forming processes, which has been described in the nineteenth century by the Russian father of soil science V. V. Dokuchaev. Podzols are soils which have an esthetically beautiful morphological appearance with a strongly leached pale albic horizon overlaying a dark, SOC- enriched, horizon as well as a bride reddish horizon colored by the Fe-oxides (Fig. 17.4).

Podzolization is a combination of leaching and translocation of SOC and cations. The humid conditions force a vertical transport of water. This process may dissolve cations like Na+, K+, Ca2+, Mg2+, and others through cation exchange and hydrolysis and leach them deep into the soil or even into the ground water and further away. When the conditions of Al-mobilization and later Fe-mobilization through decreas-ing soil pH are met, soluble organic C compounds will complex Al, Mn and later Fe, and transport it into lower horizons. The typical situation for a Podzol is that the mobilized materials are immobilized in the subsoil which may be only 2040 cm deeper than their place of origin. With respect to translocation of matter, the process is very picturesque. It does not affect large amounts of Fe and SOC. Generally, the translocated SOC is 0.61.2 kg C m2, which means 612 Mg C ha1, while the Fe loss is often only 200400 g Fe m2, which is 24 Mg Fe ha1 (Stahr 1979).

At the scale of the horizons, the process is prepared by leaching of basic cations from soil solution and from exchange complexes. The pH drops down below pH 5. Then organic compounds which are water soluble and leached out from the litter above the mineral soil are forming chelates first with Al-ions. If the pH drops below 4, it additionally mobilizes Fe-oxides. The acidification is caused first by carbonic acid (H2CO3) later by organic acids like oxalic, citric or acetic acid, and finally by strong mineral acids like nitric or sulfuric acid. The ions of Fe and Al are liberated

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by acid weathering, then bound in organic complexes and finally leached down into the sub-soil. There are multiple reasons for re-precipitation of these compounds. There is either polymerization of SOC in the subsoil, or simply the deceleration of the water fluxes, over-saturation of the organic complexes with Al and Fe or miner-alization of OM by fungi. One of the interesting side-effects of podzolization is the non-stoichiometric transport. Some mobilized compounds are not immobilized in the spodic subsoil horizon but transported laterally below the soil profile or verti-cally out of the catchment (Sommer et al. 2001). As the Podzols are generally very poor and unfertile soils, they do not promote much turnover of organic material. Therefore, even a litter production of only 1 Mg ha-1 will gradually lead to an enrich-ment of SOC, especially when the litter is not readily decomposable. Therefore, the turnover time of SOC in Podzols is higher than in other mineral soils and may be up to 2,000 years (Stahr 1979).

17.4 Soil Organic Carbon and Soil Fertility

Soil fertility refers to the amount of nutrients in the soil which is sufficient to sup-port plant life (Derek and Bogs 2009). SOC alone does not provide any essential nutrient to crops. However, since organic C is closely linked to organic compounds

containing either N or P, there are close relationships between, e.g., SOC and soil organic N over a wide range of soils. This relationship is particularly valid for cul-tivated soils where the C:N ratio in the SOM is generally in a narrow and stable range between 10 and 15. Thus, large stock of SOC is equivalent to large amounts of organic N in the soil and hence high N availability to crops, under the assumption that N mineralisation rates are similar. In fact, at the global scale, well drained min-eral soils with highest amounts of SOC storage tend to be the food baskets for quite a number of countries (Fig. 17.5). The corn (Zea mays L.) belt in Northern America coincides with regions where Chernozems and Phaeozems predominate. The same is true for the major wheat producing region in Argentina (Gran Chaco), as well as for the Chernozem belt stretching from Eastern Europe to Central Asia. The crop-land in these regions is managed using large amounts of inputs especially fertilizers and in some cases water for irrigation, which may reduce the effect of N supply from soil organic N to the crops. Nevertheless, Bauer and Black (1994) report from long-term fertilizer experiments in the Great Plains (US) with 55125 kg N ha1 year1 that 1 Mg of SOC increased wheat (Triticum aestivum L.) dry matter produc-tion by 35 kg ha1. On the other hand, in low input fallow systems of Western Africa with less than 5 kg N ha1 year1, Gaiser (1993) observed an increase of 320440 kg DM ha1 by 1 Mg of SOC with comparable total SOC stocks. The tenfold impact on biomass production in West African cropping systems cannot be explained exclusively by the higher temperature or the physiological differences between wheat and maize. This illustrates the importance of the total amount of SOC on soil

fertility and crop productivity especially in low input systems, because they almost entirely depend on SOM as a source of available N for the crops.

However, besides the total amount of SOM its composition is also relevant for the amount of mineral N to be released during the cropping season. For organic amendments to the soil the C:N ratio is considered to be the major determinant regarding its effect on the subsequent crop. If C:N ratio of applied organic materials is above 2030, there is a high risk of N immobilization following the first months after its application, causing reductions in N availability and yield (Parnas 1975). High amounts of organic manure may cause a change in SOM composition which results in changing N availability. In this case, total SOC is disconnected from N availability and hence its correlation with crop yields as shown in Fig. 17.6.

With almost the same amount of SOC, the soil amended with prunings of a legu-minous tree (Senna siamea) produces more than twice as much maize dry matter than the one amended with maize residues. However, the light fraction of the SOC, which is defined as the floating organic material with a grain size fraction between 0.5 and 2 mm, is clearly correlated with maize dry matter production. This is sup-ported by the fact that N mineralisation in these soils was closely linked to the amount of N in the light fraction (Fig. 17.7).

17.5 Soil Organic Carbon and Supporting Ecosystem Services

Soil formation and nutrient cycling are recognized as supporting ecosystem services Millennium Ecosystem Assessment (2005). Supporting ecosystem services are not directly affecting humans and the function of ecosystems but act through all other ESs because the other ESs (provisioning, regulating and cultural) depend partly or

0

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Fig. 17.6 Comparison between total soil organic carbon, the light fraction of soil organic carbon and maize dry matter production in a tropical Acrisol (based on data of Gaiser 1993)

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entirely on the supporting ESs. In the previous sections, we have shown that SOC is an important driver in soil formation and nutrient cycling and thus, SOC plays a vital role in this supporting ecosystem service. Many soil types with their inherent properties are related to SOC accumulation like the Histosols. In the global C cycle, these soils act as C sinks, as long as they are not drained or their natural vegetation is not converted into arable land. In addition, Histosols are typical wetland soils and recognized hotspots of biodiversity. The same is true for Podzols, whose genesis strongly depends on the presence and composition of SOC. These soils with their specific properties are unique habitats and are rich in distinct above-ground and below-ground faunal and floral species, thus supporting ecosystem biodiversity. Other well-drained soils with SOC accumulation (Chernozems, Kastanozems) con-stitute the national breadbaskets of many countries and they support the provision-ing ESs of specific biomes. Finally, it has been shown that SOC is an essential driver in the reduction of a series of soil compounds like nitrate, iron or manganese oxides, thus influencing the regulatory ESs in the global nitrogen and iron cycle and favoring the emission of green house gases from partly or temporarily anaerobic soils.

17.6 Conclusions

In conclusion, the interactions between SOC, soil formation and soil fertility may be summarized as follows:

The factors of soil formation determine the intensity of soil formation and the amount of SOC stored in a soil

SOC, in turn, influences many soil forming processes, thus leading to unique habitats for organisms both above- and belowground

SOC is strongly influencing soil fertility in well-drained soils

y = 13.3x + 0.13R2 = 0.84***

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The effect of SOC on crop productivity is mainly due to improved N supply to crops. However, the magnitude of the soil C effect depends on (i) the input intensity of the cropping system and (ii) the composition (or quality) of the SOM pools.

Acknowledgements We are grateful to Andreas Lehmann and Yakov Kuzyakov for providing soil profile pictures in Figs. 17.2 and 17.4.